Biogeosciences, 8, 505–513, 2011
www.biogeosciences.net/8/505/2011/
doi:10.5194/bg-8-505-2011
© Author(s) 2011. CC Attribution 3.0 License.
Biogeosciences
Changes in ocean circulation and carbon storage are decoupled
from air-sea CO2 fluxes
I. Marinov1,2 and A. Gnanadesikan3,4
1 Department
of Earth and Environmental Science, University of Pennsylvania, 240 S. 33rd Street,
Hayden Hall 153, Philadelphia, PA 19104, USA
2 Department of Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution,
266 Woods Hole Road, Woods Hole, MA 02543, USA
3 Geophysical Fluid Dynamics Lab (NOAA), 201 Forrestal Road, Princeton, NJ 08540, USA
4 Department of Earth and Planetary Sciences, Johns Hopkins Univ., Baltimore, MD, USA
Received: 6 October 2010 – Published in Biogeosciences Discuss.: 1 November 2010
Revised: 28 January 2011 – Accepted: 7 February 2011 – Published: 25 February 2011
Abstract. The spatial distribution of the air-sea flux
of carbon dioxide is a poor indicator of the underlying
ocean circulation and of ocean carbon storage. The weak
dependence on circulation arises because mixing-driven
changes in solubility-driven and biologically-driven air-sea
fluxes largely cancel out. This cancellation occurs because
mixing driven increases in the poleward residual mean circulation result in more transport of both remineralized nutrients
and heat from low to high latitudes. By contrast, increasing
vertical mixing decreases the storage associated with both
the biological and solubility pumps, as it decreases remineralized carbon storage in the deep ocean and warms the ocean
as a whole.
1
Introduction
The ocean carbon cycle plays a critical role in setting today’s
climate by determining the partitioning of carbon between
the atmosphere and ocean. Two natural ways of characterizing this cycle are the meridional transport and the vertical
profile of carbon within the ocean. The north-south gradient
of the meridional transport of carbon is primarily balanced
by air-sea fluxes of carbon dioxide, leading to the obvious
question of whether measuring such fluxes (e.g., Takahashi
et al., 2009) would provide important constraints on ocean
circulation or carbon storage. Using a suite of ocean circulation models with different parameterizations of interior mixCorrespondence to: I. Marinov
([email protected])
ing, we show that the air-sea CO2 flux is insensitive to the
circulation, in contrast to the storage of carbon by the ocean.
We analyze this contrasting behavior in terms of the
oceanic solubility and biological pumps of carbon. Cold
high latitude surface waters hold more dissolved inorganic
carbon (DIC) than warm, low-latitude waters. The solubility
pump is the result of processes whereby these cold waters
sink, injecting CO2 into the deep ocean (Volk and Hoffert,
1985). The biological pump is associated with the production of planktonic organic matter at the ocean surface and the
transport and subsequent decomposition of biogenic material
throughout the water column, resulting in a further enhancement of deep carbon concentrations.
The extent to which changes in ocean ventilation impact
the natural carbon pumps and air-sea CO2 exchange has
featured prominently in recent a number of recent papers.
Changes in the rate of convection and deep water formation in the Southern Ocean have been shown to strongly
impact the oceanic biological pump and atmospheric pCO2
(e.g. Marinov et al., 2008a, b) and have been invoked
to drive glacial-interglacial changes in atmospheric pCO2
(e.g. Sarmiento and Toggweiler, 1984; Anderson et al., 2009;
Toggweiler et al., 2006; Sigman et al., 2010). In an intriguing new study Le Quere et al. (2007) suggest that changes in
Southern Ocean ventilation associated with increased Westerlies have already resulted in a loss of biologically sequestered CO2 from the deep ocean via increased Southern
Ocean upwelling and a weakening of the natural biological
pump.
In this paper, we use a suite of models in which the vertical
mixing coefficient Kv and the lateral mixing coefficient AI
are varied, producing a range of ocean circulation. We show
Published by Copernicus Publications on behalf of the European Geosciences Union.
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I. Marinov and A. Gnanadesikan: Mixing and air-sea carbon fluxes
that while these changes produce compensating responses in
the air-sea fluxes associated with the solubility and biological
pumps, they do not produce a compensation in the inventories of carbon associated with these pumps. As a result, the
total air-sea CO2 flux is decoupled both from ocean circulation and from the storage of carbon within the ocean.
2
– In the high Ai model Kv is as in the Control model, but
Ai is set to 2000 m2 s−1 .
– In the high Ai -high Kv model, Ai is 2000 m2 s−1 and Kv
is as in the high Kv model. As discussed in Gnanadesikan et al. (2002), this combination of AI and Kv produces a pycnocline depth that largely matches that in
the Control run, but a meridional overturning circulation with a substantially different structure.
Methods
Our suite of simulations was made with the Geophysical
Fluid Dynamics Laboratory Modular Ocean Model (MOM)
version 3 and has previously been used for a number of biogeochemical dynamics studies (Gnanadesikan et al., 2002,
2004; Marinov et al., 2006, 2008a, b). The biogeochemistry
component follows OCMIP-2 specifications (Najjar et al.,
2007), with new biological production calculated by restoring our limiting nutrient PO4 to observations PO4,obs at each
time step t, whenever PO4 > PO4,obs :
Jprod (x,y,z,t) = (PO4 (x,y,z,t)
−PO4,obs (x,y,z,t) /τ for z < 75 m
(1)
where τ is the biological time scale, x, y, z are model longitude, latitude and depth and t is the time. Globally integrated PO4 is set constant in all experiments. Air-sea gas
exchange of carbon dioxide and oxygen is parameterized following Wanninkhof (1992).
The use of this simple biogeochemical model means that
nutrient uptake at the surface is tightly coupled to nutrient
supply and that surface nutrient concentrations remain relatively constant when the circulation changes. This relative
constancy has proved useful in analyzing the dynamics of
ocean carbon storage, allowing us to distinguish the separate impacts of changes in surface nutrients and circulation
on ocean carbon storage (Marinov et al., 2008b). However,
it should be recognized that in the real world nutrient supply
and uptake may not be so tightly coupled. For example, if the
rate of nutrient uptake in the subpolar North Pacific is set by
the rate of aeolian iron deposition, an increase in macronutrient supply due to deep mixing would result in an increase in
surface nutrients not captured by this biogeochemical framework.
Four different implementations of MOM3 are analyzed
here.
– Our “Control” is a model with relatively low Ai and Kv .
The diapycnal mixing coefficient Kv varies hyperbolically from 0.15 cm2 s−1 at the surface to 1.3 cm2 s−1 at
5000 m, and the isopycnal mixing coefficient Ai is set
to 1000 m2 s−1 everywhere in the ocean.
– In the high Kv model, Kv varies hyperbolically from
0.6 cm2 s−1 at the surface to 1.3 cm2 s−1 at 5000 m everywhere in the ocean, while Ai is as in the Control
model.
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Wind stresses for all of these models are taken from the
Hellermann and Rosenstein (1983) product and surface
fluxes of heat and moisture are a combination of applied
fluxes from the da Silva et al. (1994) dataset and flux corrections derived by restoring the surface temperature and salinity back to observations. As with the biological model, this
means that the primary effect of changes in mixing is to
change the interior circulation rather than the surface conditions. As shown in Gnanadesikan et al. (2003) the modeled Southern Hemisphere circulation and tropical upwelling
agree reasonably well with the predictions of the theory of
Gnanadesikan (1999), suggesting that numerical diffusion
plays a relatively weak role in these simulations. This is
likely because the forcing varies relatively smoothly over
time and the advection scheme is relatively non-diffusive.
Each of these models is run in three configurations that
allow us to isolate the impacts of the biological and solubility
effects on the carbon pump. In the ABIOTIC run biology is
turned off (Murnane et al., 1999), and the alkalinity depends
only on salinity so that the resulting CO2 distribution and airsea CO2 fluxes depend only on temperature and salinity. In
the BIOTIC run we turn the solubility pump off by setting
the ocean temperature and salinity constant everywhere at
the surface in the calculation of the air-sea CO2 exchange
calculation. Finally, in the FULL configuration we turn on
biology such that the resulting CO2 distributions and fluxes
depend on both biological and solubility effects.
3
Results and discussion
Subgridscale mixing and resulting changes in ocean circulation have a major impact on the natural carbon pumps
and on the solubility and biological air-sea CO2 fluxes after full equilibration of the ocean-atmosphere carbon system
has been achieved. Figure 1 shows results from our 4 models run in the ABIOTIC, BIOTIC and FULL configurations.
The surface to deep gradients in DIC in each of these configurations are representative of the strength of the solubility
pump (Fig. 1a), the biological pump (Fig. 1d) and the full
carbon pump (Fig. 1g) at equilibrium. Figure 1 supports our
contention that the linear separation of total DIC and total
carbon pumps into components works well, at least in the
global average, since the sum of the solubility pump (Fig. 1a)
and biological pump (Fig. 1d) closely approximates the total
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I. Marinov and A. Gnanadesikan: Mixing and air-sea carbon fluxes
507
Fig. 1. Carbon pumps (µmol kg−1 ), air-sea carbon fluxes (Pg C yr−1 degree−1 ) and northward carbon transports (Pg C yr−1 ) in 4 models
with different mixing. Top panels (a–c): abiotic carbon pump, air-sea DIC flux (positive fluxes are into the ocean) and northward DIC
transport integrated over the entire watercolumn from the ABIOTIC models. Middle panels (d–f): biological carbon pump, air-sea DIC flux
and transport from the BIOTIC only models. Bottom panels (g–i): total (full) carbon pump, air-sea DIC flux and transports from the FULL
model configurations which include both biotic and solubility pumps. Fluxes and transports are zonal and annual mean averages. Pumps
represent the zonally and meridionally integrated DIC at each depth minus average surface DIC in µmol kg−1 . For each model, the total or
full carbon pump (g) is well approximated by the sum of the abiotic (a) and biological (d) carbon pumps. Starred lines in panel (d) show
remineralized carbon (calculated as remineralized nutrients times rC:P = 117); surface-to-deep gradients in remineralized carbon represent
the organic or soft tissue pump strength assuming equilibrium with the atmosphere. Differences between the starred and unstarred lined in
panel (d) are due to carbonate and disequilibrium with the atmosphere. Intermodel differences in biological carbon pumps (unstarred lines
in panel d) can be explained by intermodel differences in remineralized carbon (starred lines in panel d), as discussed in the text.
carbon pump (Fig. 1g). Figure 1 also shows the abiotic, biological and total air-sea CO2 fluxes (middle column) and
transports (right hand column) corresponding to the pumps
in the left hand column. Note that the biological pump gradient is due to a combination of soft tissue export associated
with nutrient transport to the deep (the so-called soft tissue
or organic pump) and carbonate mineral export (the carbonate pump) and is enhanced by the slow rate of air-sea gas
exchange (Toggweiler et al., 2003b). Starred lines in Fig. 1d
show the contribution to the biological pump due to the soft
tissue pump assuming surface CO2 is in equilibrium with the
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atmosphere, i.e., this is the zonally and meridionally averaged respired carbon calculated from the remineralized nutrients multiplied by a C:P stoichiometric ratio, assumed to
be a constant in all our models.
Two important results emerging from Fig. 1 are:
I. An increase in Kv decreases the ocean carbon storage
associated with both biological and solubility carbon
pumps. The two effects reinforce each other, such that
total carbon storage decreases with larger Kv .
II. An increase in Kv increases the magnitude of both the
biological and the abiotic air-sea fluxes. The two effects
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I. Marinov and A. Gnanadesikan: Mixing and air-sea carbon fluxes
Control
cold
high Kv model
warm
low remin DIC
warmer
higher remin PO4
high remin PO4
Residual circulation
Ts
Tn
Tu
Residual circulation
Ts
Northern
Sinking
cold
Deep Ocean
Tn
Tu
Northern
Sinking
colder
Deep Ocean
high Kv − Control
dTemp > 0
heat transport
dRemin PO4 > 0
abiotic transp of DIC
convection
transport of remineralized DIC
(or transport of remin PO4)
Tu
Ts
dTemp < 0
dRemin PO4 < 0
Southern Ocean
Equator
Northern Oceans
Fig. 2. Schematic showing the impact of increased diapycnal mixing on circulation and the transport of abiotic DIC and remineralized PO4
in our GCM. Changes in the abiotic DIC transport and biological DIC transport cancel each other out. The total (abiotic + biological) DIC
transport is similar in the control and high Kv models and is northward.
compensate each other, such that the total air-sea CO2
flux changes little.
The impact of changing diapycnal and isopycnal mixing on
oceanic circulation and oceanic heat transport is summarized
in Fig. 2 (see also Gnanadesikan, 1999, 2003 and Marinov
et al., 2008b). Increasing vertical mixing Kv affects oceanic
circulation in 2 major ways:
A. Changing the mean hydrographic structure. This involves increasing the relative contribution of the Southern Ocean waters to the deep ocean via increased Southern Ocean convection and AABW formation (blue arrow in Fig. 2) as well as increasing the depth of the
tropical pycnocline.
B. Changing the residual mean circulation (TS in Fig. 2).
In detail, increasing Kv increases the conversion from
dense to light water in the tropics and also increases
the thermocline depth as more heat diffuses downwards.
This in turn increases the southward eddy transport and
the Northern Hemisphere overturning Tn and thus increases poleward transport of light waters and heat in
both hemispheres.
We will invoke mechanism A to explain most of the observed
carbon pump differences among models (Result (I) above),
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while arguing that mechanism (B) dominates the air-sea flux
differences among models (Result (II) above).
We begin by looking at the impact of vertical mixing on
ocean carbon storage associated with the solubility pump.
As seen in Fig. 3a, increasing vertical mixing produces an
increase in the importance of Southern Ocean source waters causing temperatures to decrease along the AABW pathway and in the deep ocean below 2500 m, and enhancing
the storage of carbon in these waters. However, in low latitudes higher mixing also results in deeper penetration of
surface heat, increasing temperatures above 2500 m depth.
This warming translates into less net storage of abiotic DIC
in the ocean as a whole (Fig. 3a, c), and thus a less efficient solubility pump in the high Kv compared to the control model (Fig. 1a). The three dimensional pattern of abiotic
DIC change is well correlated with the temperature change,
with a slope of −8.5 µM DIC K−1 .
We note that temperature changes alone give an indication
of the strength of the solubility pump assuming infinitely fast
gas exchange, i.e., surface CO2 in equilibrium with the atmosphere. As discussed in Toggweiler et al. (2003a), slow gas
exchange of CO2 makes the solubility pump less efficient at
storing carbon. Slow gas exchange slightly enhances the response of the solubility pump to changes in mixing, so that a
more convectively ventilated ocean tends to have less carbon
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I. Marinov and A. Gnanadesikan: Mixing and air-sea carbon fluxes
509
Fig. 3. Zonally averaged annual mean differences between the high Kv and control models: (a) temperature (◦ C), (b) remineralized DIC
(µmol kg−1 ) calculated as remineralized PO4 (model tracer) times 117. (c) Salinity normalized DIC (µmol kg−1 ) from the ABIOTIC run,
(d) biological DIC calculated from BIOTIC runs (µmol kg−1 ). Small differences between b and d are due to intermodel differences in air-sea
surface CO2 disequilibrium. Correlations (corr) and regression coefficients (r) with the temperature difference in panel (a) are shown in each
case. Globally, remineralized DIC is well correlated to temperature (corr = 0.61). Abiotic DIC differences are due primarily to intermodel
differences in temperature (corr = −0.91), and to a lesser extent to intermodel salinity differences.
than would be expected from the temperature change alone
(notice the difference between the zero lines in Fig. 3a and c).
Turning next to the ocean carbon storage associated with
biology, we begin by reviewing the nutrient dynamics that
underlie this storage. Everywhere in the ocean our biologically limiting nutrient is the sum of a preformed and a remineralized component. Preformed nutrients are the nutrients
that sink into the ocean interior without having been utilized
in photosynthetic production at the ocean surface; they signal
an inefficient biological pump (Ito and Follows, 2005; Marinov et al., 2006). Remineralized nutrients are those respired
throughout the water column and are stoichiometrically associated with remineralized carbon. The deep ocean is ventilated by the Southern Ocean and the NADW. The Southern
Ocean has relatively higher preformed nutrients (i.e., a more
inefficient biological pump) and lower sea surface temperature (SST).
Increasing Kv makes the deep ocean look more like the
Southern Ocean, decreasing preformed nutrients throughout
the deep ocean. The consequences are a “leakier” or less efficient biological carbon pump in the high Kv model compared
to the control LL model (see also Marinov et al., 2008b). For
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a constant total phosphate, an increase in preformed nutrient will correspond to a decrease in remineralized nutrient
(Fig. 3b) and carbon (Fig. 3d) along the AABW pathway and
in the deep ocean. Circulation-induced changes in remineralized nutrients and carbon explain intermodel differences
in the total biological pump (compare starred and unstarred
lines in Fig. 1d).
Taking the two pumps (solubility and biological) together,
lower total remineralized inventory and overall warmer temperatures mean that the high vertical mixing (high Kv ) model
stores less carbon than the control model in the ocean. Overall, increased influence of the Southern Ocean on the deep
ocean and a deeper pycnocline (Mechanism A) result in decreases in the efficiency of both biological and solubility
pumps.
But what are the implications for air-sea CO2 fluxes and
the meridional transport of carbon in the ocean? And how are
air-sea CO2 fluxes connected to the carbon pumps? Again,
we consider the effects of changing mixing on the abiotic,
biotic and net fluxes.
The abiotic air-sea CO2 flux follows to first order the inverse of the thermal flux defined by Keeling et al. (1993) as:
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I. Marinov and A. Gnanadesikan: Mixing and air-sea carbon fluxes
Fig. 4. (a) (−1) · thermal flux in Pg C yr−1 degree−1 . (b) Southward heat transport (PW) in four different models with different mixing. All
terms are annual and zonal averages. Thermal fluxes primarily reflect variations in heat fluxes and are calculated from Eq. (2) in the text,
using a surface Revelle factor R = 14−0.2·T and assuming 1/pCO2 ·dpCO2eq /dT = 0.0423◦ C−1 (Takahashi et al., 1993). Thermal fluxes
drive the abiotic DIC fluxes in Fig. 1b. Southward heat transports drive the abiotic northward DIC transports in Fig. 1c.
Q dDICeq
Q ∂DIC dpCO2
·
·
=
·
Cp
dT
Cp ∂pCO2
dT
Q
DIC dpCO2
=
·
·
Cp · R pCO2
dT
F=
(2)
where Q is the heat flux from the atmosphere to the ocean,
Cp is the heat capacity of water, T is temperature, and R
is the Revelle or buffer factor which shows the sensitivity
of pCO2 to changes in DIC. This thermal flux primarily reflects variations in the heat flux (Fig. 4a) and drives the abiotic carbon fluxes in Fig. 1b, while the corresponding southward transport of heat (Fig. 4b) drives the abiotic northward
transport of carbon in Fig. 1c.
Upwelling brings cold water to the surface at the Equator where it is heated up; the decrease in solubility results in
loss of carbon to the atmosphere. As this water moves at the
surface to higher latitudes in both hemispheres (e.g., through
currents such as Gulf Stream, Kuroshio or through eddy diffusive transport) it gradually loses heat and gains CO2 ; low
latitude outgassing is offset by CO2 ingassing in high latitudes.
The biological CO2 flux reflects the preformed nutrient
distribution at the ocean surface. In high latitude regions
such as the Southern Ocean, remineralized nutrients and carbon are upwelled to the surface. Since biological uptake is
inefficient here relative to the upward supply, preformed nutrients and pCO2 at the ocean surface build up, resulting in
CO2 outgassing to the atmosphere. Technically, the rate of
biological outgassing to the atmosphere is controlled by the
efficiency of conversion of deep remineralized nutrients to
surface preformed nutrients; more surface preformed nutrients signal more inefficient biology and more biological outgassing.
At steady state, large biological uptake in the subtropics is required to counterbalance high latitude biological degassing. Biological uptake is generally efficient in the subtropics; translating into near-zero surface preformed nutrients and net uptake of carbon from the atmosphere. High
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export of organic matter at low latitudes is reflected by high
remineralized nutrients and carbon right below the suface
(with highest concentrations above 1000 m depth). From a
nutrient budgeting perspective, we see a net conversion of
preformed surface nutrients to remineralized sub-surface nutrients. This conversion controls the rate of biological carbon
ingassing into the low latitude ocean.
The net air-sea CO2 exchange reflects a competition between biological and solubility effects, with the biological
air-sea CO2 flux opposing the solubility air-sea CO2 flux
throughout most of the ocean. In the low-latitudes, where
preformed nutrients are low, the abiotic effect is much larger,
and the air-sea flux reflects primarily the solubility flux. In
the high latitudes, the two are much closer, and the solubility
effect barely wins over the biological effect such that on average high latitudes are a sink for atmospheric CO2 (Fig. 1b,
e, h). Interestingly, this is very different from the total carbon pump in the ocean (reflected by the DIC surface-to-deep
gradients), which is dominated by the biological rather than
the solubility pump (Fig. 1a, d, g).
In order to understand the impact of mixing on the airsea carbon exchange, it is easiest to think of the abiotic and
biological air-sea carbon fluxes as the horizontal gradients of
the abiotic and biological transports of carbon, respectively
(Fig. 1c, f).
The depth integrated transport of abiotic carbon is a consequence of the net poleward heat transport in each hemisphere (Fig. 4b), and scales as the poleward water mass transport times the low latitude-high latitude temperature gradient. The biological meridional transport of carbon is given to
a first order by the poleward transport of remineralized nutrients, adjusted by the appropriate stoichiometric constant, and
scales as the poleward mass transport times the low latitudehigh latitude gradient in remineralized nutrients. In conclusion, we can understand changes in carbon transports in
terms of changes in the poleward mass transport (referred to
here as the residual mean circulation or TS in Fig. 2) and
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I. Marinov and A. Gnanadesikan: Mixing and air-sea carbon fluxes
changes in the meridional gradients of temperature and remineralized nutrients.
The Southern Ocean residual mean circulation is a sum
of the northwards Ekman advection and counteracting southward eddy transport. Changes in the residual mean circulation due to increased diapycnal mixing create a pattern
in remineralized nutrients analogous to the temperature pattern, with increased values in sub-surface low latitudes and
decreased Southern Ocean and deep ocean values (Fig. 3b
and a). Increased diapycnal mixing increases the poleward
residual mean circulation, transporting poleward both higher
remineralized nutrients (and thus more biological DIC) and
warmer tropical waters (and thus more heat). Additionally,
the spatial correlation between the mixing-driven temperature change and remineralized nutrient change is 0.61. Thus,
insofar as the increase in poleward heat transport is due to
warming of the thermocline waters, there is a corresponding increase in the poleward transport of biological carbon
due to a higher concentration of remineralized nutrient in the
thermocline. The resulting increase in poleward transport
of biological DIC compensates almost exactly the increase
in equatorward solubility driven (abiotic) DIC, such that total carbon transport and resulting air-sea CO2 fluxes do not
change much with mixing (Fig. 1h, i).
Revisiting Fig. 1, we see that changing isopycnal mixing
Ai has a much smaller effect on carbon pumps and air-sea
fluxes than diapycnal mixing Kv . An increase in Ai decreases
on average the large scale overturning circulation, decreases
equatorial upwelling and inhibits the large oceanic cooling
associated with the Western Boundary Current. Therefore,
changes in the abiotic air-sea CO2 flux due to increasing Ai
can locally oppose changes in fluxes due to increasing Kv .
The compensation between different pumps occurs despite
large impacts of mixing on individual pumps. Increasing Kv
from 0.15 to 0.6 cm2 s−1 increases the poleward heat transport by 0.8 PW in the Southern Hemisphere and 0.2 PW in
the Northern Hemisphere (Gnanadesikan et al., 2003) and the
peak poleward biological carbon transport by 0.6 Pg C yr−1
in the Southern Hemisphere. Lateral diffusion has a smaller
impact, increasing poleward heat transport in the Southern
Hemisphere by up to 0.4 PW and reducing it in the Northern Hemisphere, with correspondingly smaller impacts on
the carbon transport.
A final point which we wish to highlight is the dominant
role of the Southern Hemisphere. For example, the changes
in heat transports are 2–4 times as strong in the Southern
compared to the Northern Hemisphere. Similarly, the sensitivity of both abiotic and biological fluxes to changes in subgridscale mixing or wind intensity is highest in the Southern Ocean. The dominance of Southern Hemisphere impacts is also found for changes in wind intensity (not shown
here). This is also consistent with the results of Sarmiento et
al. (1998) who find large sensitivity in the Southern Ocean
biological and abiotic fluxes to changes in stratification following climate change. It is thus not surprising that when
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511
Watson and Orr (2003) analyzed a set of models with identical (OCMIP) biogeochemistry but different model physics
and different boundary conditions, they found the largest disagreement between biological, solubility and anthropogenic
CO2 air-sea fluxes in the Southern Ocean.
4
Conclusions
We have shown that increasing mixing decreases both the
solubility and the biological carbon pumps and carbon storage in the ocean but has little impact on the total air-sea
CO2 flux. The difference arises because different parts of the
ocean circulation are responsible for setting the magnitude
of the carbon pumps and fluxes. AABW formation plays a
critical role in setting the inventory of carbon and the magnitude of the net carbon pumps. By contrast, air-sea CO2
fluxes depend on the transport of warm waters with higher
remineralized nutrients from low to high latitudes. Changes
in the solubility CO2 flux with vertical mixing are explained
by underlying changes in temperature and meridional heat
transport, while changes in the biological CO2 flux with vertical mixing mirror changes in remineralized nutrients and
their meridional transport. As a result, we observe a strong
compensation between changes in the biological and solubility CO2 fluxes and transports such that the total flux and
the corresponding total meridional transport remain largely
unchanged (Fig. 1h, i).
To some extent the compensation mechanism is fortuitous
because the present sizes of the remineralized pool of nutrients and the current temperature gradient are such that
mixing-induced changes in the meridional gradients of nutrient and temperature have comparable and opposite impacts
on carbon. If the compensation mechanism were to hold true
in the real ocean, potential changes in circulation due to climate change, while strongly affecting the individual pump
strength and the natural carbon sequestration in the deep
ocean, would not have an observable impact on the total CO2
fluxes and the meridional transport of natural carbon. However, changes that affected only the biological pump (such
as changing iron supply to the surface ocean) would result
in air-sea CO2 fluxes that did depend strongly on circulation. This can be seen by recognizing that if nutrients were
depleted to zero everywhere at the ocean surface (following
iron fertilization) the biological air-sea flux of carbon would
vanish, leaving only the solubility-driven fluxes which (as
shown in Fig. 1b and c) do differ significantly between different circulation models.
Research has suggested that SST differences between
GCMs forced with different winds are amplified in coupled
simulations compared to ocean only models such as ours
where atmospheric temperature is held fixed (Gnanadesikan
and Anderson, 2009). We might therefore expect an even
larger response of abiotic and biological CO2 fluxes and carbon pumps to changes in winds or mixing in coupled models.
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The existence of a compensation mechanism between solubility and biological carbon fluxes might thus be even more
important and needs to be further tested in such coupled setups.
Acknowledgements. I. Marinov was supported by NOAA grant
NA10OAR4310092. We thank Jorge Sarmiento and Robbie Toggweiler for useful discussions. We also thank Rick Slater and
Roman Shor for technical support.
Edited by: V. Brovkin
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